Download Mantle-driven deformation of orogenic zones and clutch tectonics

Geological Society, London, Special Publications
Mantle-driven deformation of orogenic zones and clutch tectonics
Basil Tikoff, Ray Russo, Christian Teyssier and Andréa Tommasi
Geological Society, London, Special Publications 2004; v. 227; p. 41-64
doi:10.1144/GSL.SP.2004.227.01.03
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Mantle-driven deformation of orogenic zones and clutch tectonics
BASIL T I K O F F 1, R A Y RUSSO 2, C H R I S T I A N T E Y S S I E R 3 & ANDRI~A T O M M A S I 4
1Department of Geology and Geophysics, University of Wisconsin, Madison,
WI 53706, USA (e-mail: basil @geology, wisc. edu)
2Department of Geological Sciences, Northwestern University, Evanston, IL 60208, USA
3Department of Geology and Geophysics, University of Minnesota,
Minneapolis, MN 55455, USA
4Laboratoire de Tectonophysique, ISTEEM, CNRS/Universitd de Montpellier II,
34095 Montpellier cedex 5, France
Abstract: Compatible deformation between the upper crust and upper mantle is documen-
ted for a variety of ancient and neotectonic settings, suggesting that these lithospheric layers
are coupled. Areas of neotectonic deformation are also characterized by high seismic attenuation, indicating that the uppermost mantle is rheologically weak and flowing in these
regions. The flow of the mantle, both lithospheric and asthenospheric, potentially drives
deformation in continental orogenic zones. Three-dimensional models, controlled by
bottom-driven mantle flow, are proposed for obliquely convergent, transcurrent and obliquely divergent plate margins. Our analysis indicates that the absolute, and not just relative,
plate motions play a critical role in the orogenic cycle.
The Earth displays a radial compositional
stratification, which results in depth-dependent
variation in the strength of Earth materials (rheology). The theology of each of these distinct lithospheric layers was quantified through the use of
experimental deformation, and corroborated by
geological field studies and microstructure analysis. A steadily accruing set of diverse data indicates that the Earth's continental crust is coarsely
divisible into two mechanically distinct layers in
continental regions: the upper crust and lower
crust. The upper crust deforms by brittle or cataclastic mechanisms, while the lower crust
deforms by crystal-plastic mechanism (Sibson
1977; Scholz 1988). The upper mantle underlying continental crust is also mechanically
divisible into a more rigid lithospheric upper
mantle and a more mobile, fluid-like asthenospheric mantle (Brace & Kohlstedt 1980). Modifications to these rheological estimates have been
made, such as the recognition of the importance
of micas in upper crustal deformation (Imber
et al. 2001) or the effect of water in the mantle
rheology (Kohlstedt et al. 1995).
The compositional and rheological heterogeneity of the lithospheric layers was used to infer
horizontal drcollement zones in the lithosphere,
which results in decoupling of surface crustal
blocks from underlying mantle (e.g. Oldow
et al. 1990). If this scenario is true, it is an
important obstacle for understanding how lithospheric motions and deformation relate to deeper
Earth processes such as mantle convection: if
there is no mechanical connection, plate tectonics
is either independent or, at a minimum, does not
have a simple relationship to mantle convection.
This basic realization has recently emerged as
one of the most fundamental problems of
modem tectonics.
In this contribution, we synthesize geological,
microstructural and geophysical datasets and
propose a three-dimensional view of lithospheric
deformation. In a variety of settings, upper
crustal deformation (characterized by geological
mapping and geodetic information) is remarkably
similar to mantle deformation (shear-wave splitting data). Consequently, we infer that the layers
are at least partially coupled in a process we call
clutch tectonics (Tikoff et al. 2002). Alternatively
stated, detachment zones are not ubiquitous in the
ductile lower crust. Second, we use geophysical
arguments to suggest that the uppermost mantle
is composed of two distinctly different types:
(1) rigid lithospheric mantle under oceanic crust
and cratonic regions of continents; and (2) a
more fluid, 'asthenospheric-like' mantle under
mobile belts. Consequently, we use the term
'upper mantle' rather than lithospheric mantle or
asthenospheric mantle, as the rheology of the
upper mantle may depend on tectonic setting.
From: GROCOTT,J., MCCAFFREY,K. J. W., TAYLOR,G. & TIKOFF, B. (eds) 2004. VerticalCouplingand
Decoupling in the Lithosphere.Geological Society, London, Special Publications, 227, 41-64.
0305-8719/04/$15 9 The Geological Society of London 2004.
42
B. TIKOFF ETAL.
The concept of plate tectonics has never
worked particularly well in mountain belts:
plates as lithospheric columns that remain intact
through time do not exist in these settings. The
observations suggest that crustal deformation is
potentially tied in a direct way to underlying
mantle movement (e.g. convection), through a
series of subhorizontal, high-strain zones which
act to partially attach adjacent lithospheric
layers. We first present an overview of the observations, both geological and geophysical.
We then present the clutch model to explain
the observed phenomenon in the context of
mantle-driven flow. Last, we investigate the
importance of absolute plate motions for
orogenic deformation.
Observations
Lithospheric deformation
The upper crust is the best-studied lithospheric
layer in terms of deformation. Deformation is
both aseismic and seismic, and the bulk rheology
is mostly characterized by Coulomb behaviour.
The deformation mechanisms for upper crustal
deformation are cataclastic flow, pressure solution and dislocation creep. The translation and
rotation of upper crustal blocks is well demonstrated by Global Positioning System (GPS)
data. Internal strain of the upper crustal blocks
is also recorded by GPS data, although the
signal contains both transient elastic deformation
of the crust and shallow mantle (i.e. seismic
cycle) and permanent tectonic strain due to
plate motions. Standard kinematic analysis of
structural geological data allows one to constrain
three-dimensional, finite deformation. Accurate
average velocities (and strain rates) are obtainable with this technique, provided that the
timing of deformation is well bracketed (e.g.
Christensen et al. 1989; Dunlap et al. 1997).
The lack of normal seismicity indicates that
the lower crust (>15 km depth) is significantly
different theologically and/or compositionally
from the upper crust. Geological observations
on exposed lower crustal sections suggest a
bulk felsic composition for the lower crust, consistent with seismic velocity (i.e. Vp/Vs) studies
on in situ sections (Christensen & Mooney
1995). Layering in the lower crust is generally
subhorizontal to shallowly dipping, observed
both geologically and geophysically. In addition,
large sections of the exposed lower crust contain
nappes and sheath-fold structures, indicating
large amounts of plastic strain (e.g. Goscombe
1992). These structures are important as they
imply that the lower crust is pervasively
deformed.
Mantle deformation can be deduced from
shear-wave splitting datasets in a variety of orogenic zones. The result of seismic shear waves
travelling through anisotropic media is shearwave splitting, similar to optical birefringence,
in which the incoming shear wave is split in
two quasi-orthogonal waves that travel at different speeds (see reviews by Silver 1996; Savage
1999). The time lag between the two waves is
controlled by the intrinsic anisotropy of the
medium, which depends on the intensity of the
lattice preferred orientations, and by the thickness of the anisotropic layer. Typical splitting
delays in neotectonic regions are ~ 1 - 2 s. Most
of the observed shear-wave splitting for teleseismic travel paths is considered to arise in the
upper mantle, extending down to the 410 km
seismic discontinuity. While crustal rocks
contain large intrinsic anisotropies, the relative
thinness and the variability of the deformation
fabrics of the continental crust rule it out as a
major source of shear-wave splitting. Barruol &
Kern (1996) have shown that the crust typically
contributes ~ 0 . 1 - 0 . 2 s of splitting, with 0.5 s
delays in extreme cases. Moreover, kilometrescale exposures of upper mantle rocks show a
strong fabric, consistent with the interpretation
of pervasive deformation in these settings (e.g.
Christensen 2002).
Mantle fabric
Olivine LPO and shear-wave splitting. The
observed shear-wave splitting in mobile belts is
presumed to result primarily from the lattice preferred orientation (LPO) of olivine, the most
abundant mineral in the upper mantle. Olivine
displays anisotropic elastic properties (Abramson et al. 1997), with compressional seismic
waves travelling 25% faster when parallel to
the fast [100] axis relative to the slow [010]
axis. Shear-wave propagation is also affected
by this elastic anisotropy. The highest anisotropy
(c. 18%) is recorded by waves propagating in a
direction intermediate between the [100] and
the [001] crystallographic axes, and fast split
shear waves are polarized parallel to the [100]
axis of olivine. The anisotropy of olivine is relatively low compared to that of some other silicate
minerals, such as biotite. The large delay times
recorded by shear-wave splitting are due to the
thickness of the mantle sampled by the seismic
waves, rather than to the inherent anisotropy of
the material. Shear-wave splitting requires that
the olivine crystals display coherent orientations
(i.e. strong LPO) over large distances in the
MANTLE-DRIVEN DEFORMATION OF OROGENS
upper mantle (>50 km). Therefore, dislocation
creep must be active in at least the upper 250300 km of the mantle, since diffusion creep
does not produce an LPO. This inference is consistent with experimental deformation, which
suggests that dislocation creep occurs in olivine
aggregates at upper mantle temperatures and
pressures (e.g. Zhang & Karato 1995).
Seismic anisotropy, such as shear-wave splitting, results from the deformational fabric of
mantle rocks. Shear-wave splitting observations
provide information about deformation in the
mantle, allowing us to conduct structural geology
analysis of the mantle in these settings. The
relation between deformation and olivine LPO
is critical to our interpretation of seismic anisotropy data in terms of upper mantle kinematics.
Comparison between shape preferred orientation
of olivine and spinels in a large number of naturally deformed mantle rocks indicates that the
olivine [100] axis usually tends to concentrate
at low angle to the lineation. The [010] axis
either concentrates at high angles to the foliation
or forms a girdle normal to the lineation (Ben
Ismail & Mainprice 1998). This LPO suggests
a dominant activation of the (010)[100] and
(001)[100] slip systems in olivine for deformation
under upper mantle conditions (Tommasi et al.
2000). Analysis of olivine LPO evolution in
high-temperature deformation experiments, as
well as in polycrystalline plasticity models, provides some further constraints on this relation.
experiments.
Deformation experiments of both olivine single crystals and polycrystalline aggregates under high-temperature,
high-pressure conditions provide valuable constraints on: (1) how the activation of the
various slip systems, and hence the LPO, is
affected by physical parameters such as temperature, water and oxygen fugacity; and (2) how
olivine LPO reflects the flow kinematics and/or
finite strain. Flow experiments on olivine crystals, under predicted upper mantle pressure and
temperature conditions, show that the dominant
slip system is most likely (010)[100], followed
by the (001)[100] system. Under high water or
oxygen fugacity conditions, slip on the latter
system tends to dominate (Bai et al. 1991;
Mackwell et al. 1985). On the other hand, lowtemperature experiments show a dominant activation of glide on the [001] direction (Phakey
et al. 1972). Analysis of olivine LPO in naturally
deformed peridotites shows that upper mantle
deformation is generally accommodated by slip
on both (010)[100] and (001)[100] systems,
with dominance of the former. Moreover, glide
in the [001] direction is only significant during
Laboratory
43
low-temperature deformation associated with
the emplacement of mantle slabs into the crust
(Tommasi et al. 2000).
In simple shear experiments under high-T
conditions, deformation fabrics yield seismic
anisotropy with an LPO characterized by a concentration of olivine [100] axes parallel to the
elongation direction (Fig. 1). The fabrics
obtained in the various experiments deviate,
however, with respect to the orientation of
[010] axes and [001] axes. In plane-strain experiments, Zhang & Karato (1995) found point
maxima for the crystallographic [010] axes and
[001 ] axes of olivine approximately perpendicular to the shear plane and approximately perpendicular to the shear direction within the shear
plane, respectively. In contrast, Bystriky et al.
(2000) report girdles of crystallographic [010]
axes and [001] axes of olivine for torsion experiments (Fig. 1). Both types of LPO are common
in mantle xenoliths and mantle sections of ophiolites (Ben Ismail & Mainprice 1998).
We tentatively suggest that the Zhang & Karato
(1995) data represent slightly higher temperatures
(1300 ~ v. 1200 ~ than the experiments of
Bystriky et al. (2000). The LPO patterns are the
same, only the development is faster for the
Zhang & Karato experiments because of stronger
dynamic recrystallization. This inference is
broadly consistent with the results of Carter &
Av~ Lallemant (1970), in which point distributions of crystallographic axes and activation
of (001)[ 1001 occurred at the highest-temperature
conditions. A major difference, however, between
the Zhang & Karato (1995) and Bystriky et al.
(2000) experiments is the type of strain. The
experiments of Bystriky et al. (2000) are simple
shear, whereas the experiments of Zhang &
Karato (1995) contained a component of coaxial
deformation parallel to the shear direction
(reported in Zhang et al. 2000). This coaxial component, in addition to the possible formation of
extensional shear bands, will lower the angle
between the long axis of the finite strain ellipse
and the shear plane. Therefore, in both experiments, the [100] axes are parallel to the long
axis of the finite strain ellipsoid.
experiments.
The development of
LPO in olivine aggregates has also been
addressed using forward modelling. Several
approaches have been utilized, including viscoplastic self-consistent (Wenk et al. 1991; Wenk
& Tome 1999; Tommasi et al. 1999), stress equilibrium (Chastel et al. 1993) and kinematic constraint (Ribe & Yu 1991; Kaminski & Ribe 2001)
models. These LPO predictions using these
various approaches and their application to the
Numerical
44
B. TIKOFF E T A L .
Zhang & Karato, 1995
a[loo]
vertical foliation
vertical lineation
b [010]
horizontal foliation
horizontal lineation
c [OOl]
vertical foliation
horizontal lineation
Bystriky et al., 2000
a[loo]
b [01o]
c [OOl]
no splitting
vertical foliation
vertical lineation
horizontal foliation
horizontal lineation
vertical foliation
horizontal lineation
Fig. 1. Olivine microstructures produced under different deformation boundary conditions. The experiments of
Zhang & Karato (1995), potentially applicable to higher-temperature deformation, produce point maxima. The results
of Bystriky et al. (2000) indicate a point maximum for the [ 100] axis and a girdle distribution for the [010] and [001]
axes. The resultant shear-wave splitting, in either case, depends on the orientation of the LPO since SKS waves have a
near-vertical incidence. In either case, vertical foliation and horizontal lineation produces strong shear-wave splitting.
prediction of seismic properties are reviewed in
T o m m a s i et al. (2000).
In general, classical polycrystalline plasticity
models (in which deformation is a c c o m m o d a t e d
by dislocation glide only) predict that the olivine
[100] axes tend to parallel the m a x i m u m finite
elongation direction and that the [010] axis will
tend to align normal to the foliation. In more
complex three-dimensional flow types, the
olivine [100] axis tends to align with the long
axis of the finite strain ellipsoid (Tommasi et al.
1999). These results hold true specifically for
MANTLE-DRIVEN DEFORMATION OF OROGENS
transpressional and transtensional deformation.
These deformations are unusual because a
switch in lineation orientation (transpression) or
foliation orientation (transtension) can occur for
either: (1) different angles of oblique convergence/divergence; or (2) increasing strain at a
single angle of oblique convergence/divergence
(Tikoff & Greene 1997).
Analysis of olivine LPO development in polycrystalline plasticity models also highlights that
LPO intensity does not depend linearly on
finite strain (Tommasi et al. 2000). At shear
strains of 1-2, clear LPOs are already formed.
This result is in agreement with LPO evolution
in simple-shear experiments, which shows a
fast evolution up to shear strains of 1 or 2 and
then maintenance of a steady-state fabric
(Bystriky et al. 2000). Comparison between
model predictions and LPO intensity in naturally
deformed rocks also suggests that LPO only
records part of the total strain accommodated
by the rock.
Mantle viscosity: the significance o f seismic
attenuation
The degree of observed shear-wave splitting for
mantle paths attains a global average of just
over 1 s and is as high as 2.7 s (MarsonPidgeon & Savage 1997). These values imply
that, for nominal values of intrinsic anisotropy
derived from laboratory studies, coherent deformation fabrics in the upper mantle of the order
of 100-200 km thick exist (Silver 1996). The
vertical extent of this pervasive deformation
has given rise to an important debate (e.g.
Vinnik et al. 1989a, b; Silver & Chan 1991). Is
this pervasive fabric contained in a highly
viscous lithosphere, or does the shear splitting
signal arise from a plastically flowing asthenosphere? Given normal geotherms, especially for
tectonically active regions, 200 km of coherently
deformed material may include an anisotropic
asthenospheric source layer as well as the
'frozen' fabrics of the overlying lithosphere. If
the anisotropic layer is strictly asthenospheric,
then this ductile layer is potentially a zone of
weak viscous coupling, or even decoupling,
between the lithosphere and the deeper mesosphere. Seismic attenuation has the potential to
resolve this issue.
Seismic attenuation is due either to scattering
of seismic energy along the wave path or to
intrinsic attenuation of seismic energy caused
by conversion of displacive energy into thermal
energy (thermal dissipation). Normally, such loss
is higher for shear waves than for compressional
45
waves. Therefore, a comparison of the frequency
contents of S waves relative to P waves can be
made to be diagnostic of attenuation in a quantitative way. This measure is referred to as 'Q' for
quality factor. Q is a dimensionless number that
is inversely related to attenuation. For Earth
materials that are near their solidus, Q is
normally in the 100-200 range. Q for rigid lithosphere of oceanic slabs is 800 or more (e.g. Okal
& Talandier 1997). Consequently, Q measurements are useful indicators of rheology.
Attenuation along paths restricted to oceanic
lithosphere, excluding spreading ridges, is low
(high Q). In cratonic regions, presumably underlain by thick, cold, high-velocity roots, seismic
attenuation is also low. For these two areas,
oceanic lithosphere and cratonic regions, the
upper mantle is inferred to be relatively rigid
(e.g. Lay & Wallace 1995). In contrast, for
seismic travel paths that include oceanic asthenosphere or upper mantle beneath most actively
deforming regions, attenuation is high (low Q).
Seismologists typically interpret such regions to
be near solidus temperatures and rheologically
weak or even flowing. The presence of small
percentages of partial melt (1-2%) is often discussed for these regions (e.g. Roth et al. 1999).
The crustal contribution to attenuation is
minimal, due to the short path lengths for nearvertical, teleseismic rays.
The combination of shear-wave splitting
measurements and seismic attenuation observations can constrain the ultimate source of
mantle deformation. For example, one can discriminate between active mantle flow in a
rapidly deforming layer and presence of a relict
fabric in a lithospheric volume that is not
deforming actively. In both cases, we might
observe a strong, consistent shear-wave splitting
pattern across an area. In the flowing asthenosphere case, however, attenuation is high (low Q).
In contrast, attenuation will be low for the
preserved old lithosphere fabric. If both strong
lithospheric and asthenospheric fabric are
present in an area, Q measurements may still discriminate between the two given the proper set of
earthquake sources and seismic stations.
The lateral extent of long delay-time shearwave splitting measurements in active tectonic
regions suggests that the relatively narrow
plate boundaries observed on the surface in continental regions correspond to very wide plate
boundaries in the upper mantle (Russo et aL
1996; Klosko et al. 1999). Assuming that this
upper mantle material is flowing by crystalplastic processes, lateral displacements are
potentially controlled by variations in lithospheric
thickness. Convergent boundaries, by definition,
46
B. TIKOFF ET AL.
result from the juxtaposition of two different
lithospheres. Consequently, mantle flow in
these regions is affected by thickness variations
in both lithospheres.
Lithosphere/asthenosphere connections
Cratons and asthenosphere
Oceanic lithosphere contains a seismically detectable low-velocity zone inferred from dispersion of
surface waves in ocean basins. Continental lithosphere, in contrast, lacks a low-velocity zone.
Additionally, tomographic observations indicate
that Precambrian cratonal regions are underlain
by high-velocity regions extending to 300 km or
more in some cases (e.g. van der Lee & Nolet
1997). These observations imply that: either (1)
cratonic regions are coupled to large-scale
mantle circulation (and hence there is no asthenospheric shear zone at their bases); or (2) any
decoupling layer beneath cratons is much deeper
than
decoupling
beneath
ocean
basins
(~ 160 km) and is essentially different in seismic
character. Compositional differentiation whereby
a buoyant, strong residuum of mantle material
forms the roots beneath ancient continental cores
can account for the thicknesses, observed seismic
velocity and apparent unsubductibility of the
cratonic mantle lithosphere (e.g. Jordan 1975;
Carlson 1995). These roots may also explain
that asthenosphere does not develop at normal
depths beneath these regions. If 'asthenospheric'
shear zones do exist beneath cratons, the thermal
regime, pressure regime and composition are
significantly different from those of oceanic asthenosphere. Therefore, significantly different
rheology and strain rates are expected. Given the
possible viscosity increases, the effects of these
differences on rheology are difficult to constrain
quantitatively.
Insight into the deformation regime beneath
cratons can be garnered from mantle xenoliths
brought to the surface at kimberlite pipes.
Kennedy et al. (2002) note that mantle peridotites
that demonstrably come from the deep subcratonic mantle are divisible into two groups:
(1) coarse-grained protogranular peridotites, and
(2) fine-grained porphyroclastic peridotites. The
latter are pervasively deformed at higher strain
rates than the former and are only found at
lower lithospheric levels ( ~ 1 5 0 - 2 0 0 k m ) as
determined from thermobarometry. Kennedy
et al. (2002) infer that deformation recorded by
the porphyroclastic peridotites is more recent
than the last-laaown deformation (Phanerozoic)
visible at the Earth' s surface. What younger deformation can these fabrics possibly represent?
These authors argue that the porphyroclastic peridotites found in the kimberlite xenoliths are
samples of the lithosphere-asthenosphere decoupling shear zone beneath cratons, thus explaining
their high strain rate deformation (Kennedy et al.
2002). If this is correct, then shear at the base of
cratons occurs deeper and without the associated
seismic low-velocity zone found in oceanic asthenosphere. There are several potential problems
with this interpretation, including the lack of independently confirmed textural ages of the fabrics
and the cause of initial formation.
Finally, we note that the 'high' strain rate
inferred from these xenoliths may be only relatively high, in comparison to those expected
in oceanic asthenosphere. Thus, the very low
seismic attenuation typically observed in cratonic regions, and the absence of a seismically
defined low-velocity zone, could both be consistent with a deep, lower strain rate acting within a
decoupling zone beneath cratons. Moreover,
subcratonic mantle shear zones have sufficient
time to acquire such a fabric. In contrast, the
fabrics developed in oceanic asthenosphere are
limited by the temporal existence of unsubducted
young oceanic lithosphere (< 170 Ma).
Lithos_pheric v. asthenospheric fabrics
Oceanic settings. The relation between lithospheric and asthenospheric upper mantle 'deformation' fabrics is clearest in the Pacific Ocean
basin. Pn waves (Moho head waves, which are
a regional seismic phase that travels along the
Moho for most of its propagation) travelling
just below the oceanic crust exhibit strong azimuthal anisotropy (e.g. Hess 1964; Raitt et al.
1969). The fast trends for these phases align
with the mineralogical fabric imparted at spreading ridges when magma freezes while undergoing shear at the base of forming lithosphere
(Fig. 2). This shear flow aligns upper mantle
olivine [100] axes parallel to the spreading direction. In the older portions of the Pacific basin,
this alignment results in fast Pn trends parallel
to fracture zones. In contrast, a small but
growing dataset of teleseismic shear-wave splitting indicates that this anisotropy is not detected.
Instead, a strong asthenospheric signal yields fast
shear polarization directions parallel to the current
absolute plate motion (hotspots assumed fixed) of
the Pacific plate (Fig. 2; Russo & Okal 1998;
Wolfe & Silver 1998; Klosko et al. 1999).
Surface wave analyses of the ocean basins show
that upper mantle azimuthal anisotropy is generally closely related to lithospheric absolute plate
motions, particularly in the Pacific basin
(Nishimura & Forsyth 1988, 1989; Montagner &
MANTLE-DRIVEN DEFORMATION OF OROGENS
Qa
~
%
'q
I
---~--~ spreading center
plate motion
~
lithospheric shear wave splitting
< ~
asthenospheric shear wave spirting
Fig. 2. Shear-wave splitting in oceanic material, in
map view. Lithospheric fabric, frozen in from
asthenospheric flow at the spreading centre, is parallel
to the fracture traces. Asthenospheric fabrics are
parallel to the current direction of plate motion in a
fixed mantle reference frame. These two diverge if the
motion of the plate has changed, such as for oceanic
material older than 43 Ma in the Pacific region.
Tanimoto 1990, 1991). Pn anisotropy and shearwave splitting are generally not parallel in the
Pacific basin, since the fracture zones in lithosphere older than around 43 Ma do not parallel
today's absolute plate motion direction.
These data have been interpreted to signal the
operation of simple asthenospheric flow beneath
the Pacific plate (Nishimura & Forsyth 1988, 1989;
Montagner & Tanimoto 1990, 1991; Russo &
Okal 1998; Wolfe & Silver 1998; Klosko et al.
1999; Vinnik et al. 1989a, b). As the plate
moves rapidly towards western Pacific subduction zones, shear is imposed by the underlying
asthenosphere. This shear transfer is presumed
to occur in the low-velocity zone, producing
effective decoupling or weak viscous coupling
between the plate and the upper mantle below
the asthenosphere.
Fabric in the lithospheric mantle is not properly a
lithospheric deformation fabric in the sense we have
been using elsewhere in this chapter. It records a
'frozen' asthenospheric flow that formed at a spreading centre, and not lithospheric deformation. This is
distinct from the current fabric in the asthenosphere,
which formed since 43 Ma when the Pacific plate's
motion changed. Thus, asthenospheric fabric
beneath oceanic lithosphere is potentially erased
in an oceanic setting. This 'erasing' may not occur
in continental lithosphere-asthenosphere fabrics
(Vauchez & Garrido 2001).
47
Continental settings. The more complicated
horizontal layering of the continental lithosphere-asthenosphere system produces a different
response to tectonic loading. Lateral loading has
the three end-member cases of contractional,
extensional and wrench loading. Each type of
loading results in a different mechanical
response: thrusting, viscous shortening, thickening and buckling (contraction); normal faulting,
crustal boudinage and extensional necking
(extension); and strike-slip faulting and ductile
shear flow (wrench). However, the observational
basis for comparing lithospheric and asthenospheric fabrics in continental lithosphere is
lacking for two reasons. First, discrimination
between lithosphere and asthenosphere in continental settings requires more sophisticated
seismic measurements than standard travel time
inversion. Second, the heterogeneity of the continental lithosphere makes resolution of the lithosphere-asthenosphere difficult. For example,
tomography might reveal what appears to be a
cold, stiff lithosphere overlying a hotter asthenosphere, but a compositional difference could just
as well explain the two velocity anomalies.
Attenuation measurements for phases following
specific paths, such as Pn or Sn, can afford a
basis for discrimination between layers that are
rheologically strong or weak. However, interpretations are tenuous without earthquakes or
well-placed stations. Pn (Moho head) waves
appear not to propagate in geologically complex
regions, although the exact reasons are unknown.
This leads to an unsatisfying guess about the
actual rheology at depth.
Regions that are tectonically active are least
likely to have thick lithosphere (e.g. Plomerova
et al. 2002), and thus measurements of core
phase shear-wave splitting (e.g. SKS) are likely
to sample asthenospheric anisotropy only. Even
if the lithospheric signal is clearly discriminated
from the asthenospheric signal, the paths of shear
waves most commonly used for splitting analyses cross both, yielding an inherent uncertainty
in the source(s) of any splitting observed at the
surface. In a few continental regions, systematic
variability in splitting parameters was detected
and interpreted as a function of the arrival azimuth
of the shear wave at the station (Savage & Silver
1993). Such variability is most often ascribed to
the presence of two layers of different anisotropy,
each of which splits the arriving shear waves.
Interpretations of such data have so far been
restricted to assuming that the topmost layer is
the lithosphere (or crust) and the lower layer is
the asthenosphere (or possibly the mantle lithosphere) (Savage & Silver 1993). The inherent
uncertainty arises from the mismatch between
48
B. TIKOFF E T A L .
two assumed layers of anisotropy, although there
are four potentially different anisotropic layers to
consider (upper crust, lower crust, mantle lithosphere and asthenosphere). As the crustal component contributes only 0.1-0.2 s of splitting
because of its limited extent (Barruol & Kern
1996), the mantle layers should dominate the splitting. Where two-layer anisotropy is detectable,
different fabrics must be stacked vertically, which
requires strong vertical deformation gradients.
An advanced type of study involves simultaneous determinations of Pn velocity and
anisotropy (Hearn 1996, 1999; Mele et al. 1998;
Calvert et al. 2000). Pn can provide a useful
constraint on the question of lithosphereasthenosphere fabric differences. This phase preferentially samples the uppermost mantle, which
is probably lithosphere even in active tectonic
environments with thin lithospheres. Estimates
of azimuthal anisotropy from Pn can be directly
compared to core phase shear-wave splitting
measurements, which generally sample the
entire upper mantle, including the lithosphereasthenosphere package.
Furthermore, seismic investigations require a
trade-off between Pn anisotropy and lateral velocity (Mochizuki et aL 1997), compromising
fine resolution. However, careful studies show
interpretable results (Hearn 1996). Good Pn anisotropy is comparable to shear-wave splitting
results, and can potentially determine whether
the splitting is dominated by the uppermost
mantle portion of the phases' travel paths or
deeper portions thereof. If the Pn and splitting
anisotropies are similar, the vertical distribution
of anisotropy must be fairly constant. Moreover,
if mantle lithosphere and asthenosphere are both
present, the fabrics of each must be similar.
Alternatively, if the anisotropies differ, it is
likely that the lithospheric fabrics differ from
asthenospheric fabrics.
In many regions, comparison of Pn- and
splitting-derived anisotropies indicates that the
uppermost mantle fabrics are very similar to
the integrated fabrics delineated by shear-wave
splitting. The tectonic environments in which
this is true include the region around the San
Andreas fault and the western portion of the
Great Basin (Hearn 1996), the Cascades region
of western Washington state (compare Hearn
1996 and Bostock & Cassidy 1995), the central
Appenines (Margheriti et al. 1996; Amato et al.
1997; Mele et al. 1998; Hearn 1999) and the
Aegean region (compare Hearn 1999 and
Hatzfeld et al. 2000). Thus, regions dominated
by contraction, extension and wrenching all
include large regions where uppermost lithosphere
and deeper lithosphere-asthenosphere packages
have similar fabrics in the rheologically layered
continental stack. Note that the correspondence
between Pn and splitting anisotropy is not
always close, even in the small portion of the
continents that have been carefully studied.
Thus, there is the possibility for differential vertical deformation between the lithosphere and
asthenosphere. Regions that document disagreement between the two sets of observations
include the Sierra Nevada region and central
valley of California, the Idaho Batholith-Snake
River Plain region, the Rio Grande rift area
(Hearn 1996), the Tyrrhenian Sea (Mele et al.
1998; Hearn 1999) and the Rhine Graben
(Vinnik et ai. 1994; Enderle et aL 1996; Plomerova et al. 1998; Hearn 1999).
Coincidence of crustal deformation and
mantle fabric
Ancient orogens
In a variety of settings, upper crustal and mantle
deformation are remarkably similar. In all cases,
mantle deformation is constrained by shear-wave
splitting data. The deformation of the upper crust
is recorded by a variety of features. In ancient
orogens, the record of upper crustal deformation
is typically that of the orientation of structures.
The Superior Province of North America and
the Kaapvaal craton of South Africa both show
shear-wave splitting parallel to the dominant
Archaean-age structure of the region (e.g.
Silver 1996). Likewise, parallelism between
Late Proterozoic collisional fabrics in the mid
to lower crust and SKS splitting data are also
recorded in the Ribeira-Aracuai belt of SE
Brazil (Vauchez et aL 1994; James & Assumpcao
1996). Palaeozoic orogenies show similar patterns. Barruol et al. (1997) show that the shearwave splitting parallels the structural trend of
the Appalachian-Caledonian-Hercynian belt.
Consequently, for this and all the above cases,
the current shear-wave splitting data represent a
'frozen' signal in the mantle lithosphere from
orogenesis.
Vauchez et al. (1997) further suggest that the
reactivation of orogenic belts in continental
break-up is a result of this 'frozen' anisotropy
in the mantle lithosphere. Polycrystalline plasticity models show that a mechanical anisotropy,
due to olivine LPO frozen in the lithospheric
mantle since the last orogenic episode, may
induce a directional strain softening in the lithospheric mantle and control the localization of
continental break-up (Tommasi & Vauchez
2001).
MANTLE-DRIVEN DEFORMATION OF OROGENS
Neotectonic orogens
Neotectonic zones allow more quantitative conclusions to be drawn on the relation between the
development of mantle fabric and crustal deformation. A wide range of neotectonic areas have
been studied, ranging from transcurrent domains
(Trinidad) to convergent ones (Himalayas). The
San Andreas system and New Zealand represent
intermediate situations, characterized by increasing amounts of convergent motion in obliquely
convergent settings (Fig. 3).
Despite the wide range of tectonic settings,
there are clear similarities in all of the above
examples. Regardless of the angle of relative
plate motion, the fast polarization direction of
the shear-wave splitting is often at a small
angle to the plate boundary. This is particularly
clear in the obliquely convergent boundary in
New Zealand (Fig. 3; Klosko et al. 1999;
Molnar et al. 1999). The stations displaying the
highest delay times show a fast shear wave
polarized at ~ 1 0 - 1 5 degrees from the trace of
the Alpine fault. As one moves away from the
plate boundary, the magnitude of the time
delay decreases and the directions are oriented
at higher angles to the Alpine fault. This observation indicates that some rotation in the orientation of LPO occurs as a result of finite strain
gradients. Patterns of shear-wave splitting show
ubiquitous rotation of splitting directions near
strike-slip faults (San Andreas fault, California;
E1 Pilar fault, Venezuela) in obliquely convergent
systems. This observation suggests that zones of
strain localization in the upper crust correspond
to zones of strain localization in the upper
mantle (e.g. Molnar 1992). In the Himalayas,
fast shear-wave polarizations are only parallel
to the trend of the belt close to the major
strike-slip faults (Fig. 4; Lave et al. 1996). The
general orientation of shear-wave splitting
subparallel to the plate boundaries suggests:
(1) rotation of the finite strain ellipse to an orientation parallel to the boundary; or (2) lateral flow
along the margin (plane transpression of
Tommasi et al. 1999).
Interpretation o f the data
Holt (2000), using geodetic data from Tibet, calculated maximum shear strain rate planes. One of
these planes is typically subparallel to the observed
shear-wave splitting. Three different approaches
were tried in South Island, New Zealand. Moore
et al. (2002) utilized the orientation of the present
velocity field and assumed that the deformation
had accumulated for ~6.5 Ma with this velocity
field. The resultant deformation fits well with the
49
observed shear-wave splitting. Little et aL (2002)
calculate surface deformation using deformed,
crustal-scale markers, and also find a very good
correlation to inferred mantle deformation. In
general, the surface deformation and shearwave splitting data are consistent in extent and
magnitude. It is not yet clear how to interpret
this coincidence. There are at least two possibilities: (1) the mantle and the crust are coupled
(McKenzie & Jackson 1983; Molnar 1992;
Teyssier & Tikoff 1998); or (2) similar boundary
conditions apply to the upper crust and upper
mantle (Holt 2000). A third possibility, simple
asthenospheric flow below the continents (e.g.
Silver 1996), does not apply to orogenic zones.
An exception to the last possibility is the proposed flow of asthenosphere under the Basin
and Range of the western United States (e.g.
Silver & Holt 2002).
The possibility that both lithospheric layers
respond independently to side-driven boundary
conditions (e.g. Holt 2000) is unlikely for a
variety of reasons. First, the upper crust, as an
individual lithospheric layer, is unlikely to transmit the stresses from the plate boundary to the
interior of Asia. This is particularly true for the
Himalayas given the existence of multiple lithospheric-scale suture zones bounding the collisional terranes, which would act as zones of
strain localization. Second, Giorgis et al. (2004)
document the coincidence in upper crustal deformation and mantle deformation associated with
vertical-axis rotations in dominantly transcurrent
environments. They document that bottomdriven models, compared to side-driven models,
better describe the amount of vertical-axis
rotation relative to the orientation of shearwave splitting. The most significant, however,
is the nature of mechanical instabilities within
independently deforming layers. The thin and
laterally extensive geometry of the lithospheric
plates allows for only a few styles of deformation
in response to a horizontal end load, such as
buckling or boudinage (e.g. Burg et al. 1994
for the Himalayas). These types of mechanical
instabilities cause localization of strain, which
leads to discontinuous deformation. However,
geodetic surveys of most orogenic zones indicate
that upper crustal flow is broadly continuous (e.g.
Beavan & Haines 2001 for the South Island of
New Zealand). Consequently, it is unlikely that
the upper crust deforms independently of the
rest of the lithosphere.
An equally viable alternative is that the layers
are in mechanical communication. This requires
that the lower crust, and other presumed weak
lithospheric horizons in the lithosphere, act to
partially couple the upper crust and lithospheric
50
B. TIKOFF E T A L .
Fig. 3. The relation of absolute plate motion (APM), relative plate motion (RPM) and shear-wave splitting (dark
lines) for four neotectonic areas. See text for details.
MANTLE-DRIVEN DEFORMATION OF OROGENS
51
Fig. 4. Cartoon of lithospheric deformation in Tibet. Dark lines indicate the orientation and magnitude (length) of
shear-wave splitting measurements. The short, wavy lines indicate zones of ductile shearing. Crustal deformation is
characterized by normal faulting and strike-slip faulting, in addition to contractional structures. Mantle deformation,
inferred from shear-wave splitting, indicates components of sinistral simple shear, concentrated near strike-slip
faults, and coaxial extrusion (left side of diagram - shown with kinematic axes a, b and c). The coincidence of crustal
deformation, inferred from geodetic motion in addition to geological structures, with the mantle deformation
indicates that the lithospheric layers are coupled.
mantle layers. We explore this option in the next
section (Clutch Tectonics).
shear-wave splitting. Because shear-wave splitting
results from a fabric, it is critical to incorporate the
geological history of the plate boundary.
A m a j o r limitation: strain history
Little et al. (2002) suggest that the geological
history is potentially recorded in seismic anisotropy. Using two markers that were offset by the
Alpine fault in South Island of New Zealand, the
Junction magnetic anomaly and the Moonlight
fold belt, the finite strain was modelled for upper
crustal deformation assuming a transpressional
deformation model. The calculated upper crustal
deformation is in very good agreement with the
shear-wave splitting. While this approach has the
disadvantage of the assumption of originally
linear markers, it has the advantage of addressing
the geological history. Little et al. (2002) demonstrated that the upper crustal deformation was
probably not constant with time, consistent with
plate motion reconstructions. Therefore, it is
potentially incorrect to determine the currently
active deformation (e.g. geodetic data) and then
integrate the data over some time to calculate the
Clutch tectonics
We hypothesize that deformation in the crust and
mantle are similar because these lithospheric
layers are mechanically connected. Therefore,
either displacements initiate in the upper crust
and move downward (top-driven system) or displacements initiate in the mantle and move
upward (bottom-driven) (Fig. 5). The zone that
connects the top to the bottom is called a clutch
zone, by analogy to the clutch of an automobile
(Tikoff et al. 2002). A clutch allows mechanical
continuity between the independently operating
wheels and engine. Clutch zones, when present,
have a different shear sense for top-driven or
bottom-driven systems (Fig. 5c v. 5d). While
the concept of top-driven systems was investigated for extensional deformation (e.g. Axen
et al. 1998), bottom-driven systems must be
the general rule in orogenic zones if tectonic
52
B. TIKOFF ETAL.
Original position
KEY
shear
sense
indicator
l
Lithosphere
Asthenosphere
a,
-w
Decollement
I Lith~
i
]
I
Asthenosphere
Top moves,
bottomdoes
notmove
Top-driven
I Lith~
T
I /Top movesfaster
~'N~(~.
ere~ thanbOttOmand
Asthenosph "drags"it along
Bottom-driven
Lithosphere I
fA ~
f
d.
sthenosphere
Bottom-driven
J~l[~" Lithosphe~
e. , #Asth~en~
f
l
Bottommovesfaster
thantopand"drags"
it along
l Bottommovesfaster
thantopand"drags"
it along
Fig. 5. Vertical, cross-sectional views of the relation between crustal and mantle deformation. The indicator
represents an imaginary, initially vertical line that changes orientation with deformation (a loose line). The crust
moves the same amount in all cases, but the relation between the upper crust and the mantle changes. Detachments (b)
cause no fabric and are not consistent with the coincidence of crust and mantle deformation. Top-driven systems (c)
produce the opposite sense of shear from bottom-driven systems (d,e). The lower crust can be involved (e), or not (d),
in a bottom-driven system. Bottom-driven systems are expected to be the norm in orogenic belts. Modified from
Tikoff et aL (2002).
m o v e m e n t s ultimately result from deep-mantle
processes. Bottom-driven systems do not require
particular wavelengths of deformation, as the
deformation does not proceed by amplification
of instabilities. Rather, deformation is semicontinuous, an observation which is matched in
many orogenic zones (e.g. Beavan & Haines
2001).
Top-driven systems
The Pacific Ocean l i t h o s p h e r e - a s t h e n o s p h e r e
(Fig. 2) and other ocean basins are potentially a
good example of a top-driven tectonic system.
In this case, the m o v e m e n t of the oceanic lithosphere may be governed by boundary forces
affecting the edges of the plates (ridge push or
slab pull), which affects the flow of the asthenosphere. Available seismic anisotropy data imply
a velocity gradient b e t w e e n the plate and the
mesosphere, and it is assumed that the higherviscosity mesosphere moves more slowly.
Regardless, as the lithosphere is pulled toward
a subduction zone, the gradual variation in viscosity that causes the rheological layering of
the lithosphere and asthenosphere will lead to a
MANTLE-DRIVEN DEFORMATION OF OROGENS
wide zone of decoupling. Note that this topdriven system does not deform the oceanic
crust away from the ridge. This rigidity probably
results from the homogeneous nature of oceanic
crust compared to continental crust.
Bottom-driven systems
The coincidence of upper crustal and upper
mantle deformation in continental settings
suggests that the processes in these two layers
are coupled. Movement must ultimately result
from mantle convection driven by thermal dissipation, and thus the likely outcome of coupled
mantle and crustal deformation is that orogenic
systems are bottom-driven (e.g. Molnar 1992).
The formation of fabrics in the lower crust is
also consistent with a notion of partial attachment
(Tikoff et al. 2002). Even complete attachments
result in the formation of different fabric, as the
lithospheric layers have differing rheologies.
Complete d6collements occur in Nature and
require zones of mechanical decoupling, which
are generally discrete (e.g. salt horizons in a
thrust system). The ubiquitous occurrence of
originally subhorizontal, lower crustal layering
and structural indications of high strain (intrafolial folds, sheath folds) suggests mechanical
communication between the upper crust and
upper mantle. Consequently, not only is
coupled surficial (upper crustal) and mantle
deformation explained by clutch tectonics, but
so is the formation of lower crustal layering
(Tikoff et al. 2002).
The major question of what causes the flow
patterns in orogenic zones remains. If one of
the lithospheres involved in plate interaction is
relatively strong, such as an oceanic or strong
cratonal region, deformation will preferentially
occur in the weaker region, which will eventually
become a deformation zone. The effect of lithospheric roots (from cratons, magmatic arcs, etc.)
is therefore to channel asthenospheric mantle
flow. The three-dimensional flow in the mantle
within these deformation zones is thus controlled
by relative plate movement. This movement, in
turn, causes pressure gradients in the mantle and
lithospheric thickness variations. Finally, part of
the simple shear component parallel to the plate
margin is probably an intrinsic part of the global
(convective) mantle flow. It is clearly obtained
in convection models, provided that the plate
boundaries are characterized by a very efficient
strain softening.
Implications for tectonics
Based on surficial crustal movements, interpretation of lithospheric and asthenospheric mantle
53
fabric, and inferences of lower crustal deformation, we propose three-dimensional models
for obliquely convergent, transcurrent and obliquely divergent boundaries. These are shown in
block diagrams in Figure 6. We first discuss the
strain patterns (e.g. Fossen & Tikoff 1993;
Fossen et al. 1994) and the resultant LPO (e.g.
Tommasi et al. 1999) that we would expect if
the deformation were homogeneous. We then
compare these predictions to the observations
of crustal deformation and shear-wave splitting.
We wish to emphasize that these models are
not strictly based on lithospheric strength profiles. Although these strength profiles are the
basis for many tectonic models, they are extrapolations of laboratory data that may or may
not be useful guides to rock behaviour at geological conditions. While strength profiles are widely
utilized, our approach is to relax these criteria to
put together an alternative, empirically based
model. Additionally, we do not explicitly incorporate other models, such as lower crustal
channel flow (e.g. Royden et al. 1997), which
are derived principally from the strength profiles.
T r a n s c u r r e n t b o u n d a r i e s (Fig. 6a)
Wrench deformation makes a very simple prediction about the formation of fabrics. The
material is essentially deformed within a simple
shear zone, with the shear zone oriented vertically
and parallel to the plate margin. In this setting,
plane strain fabrics are expected to develop.
Foliation is vertical and lineation is horizontal
(foliation is defined by the maximum and intermediate axes of finite strain; lineation by the
maximum axis of finite strain). Although both
fabrics are initially oriented at a 45 ~ to the
shear plane, they rotate into parallelism with
increasing finite strain (e.g. Lin & Williams
1992). The fabric in the mantle that corresponds
to this style of deformation depends on the exact
process of formation of the LPO (Fig. 1; e.g.
point distributions v. girdle patterns). Regardless
of the model chosen, strong shear-wave splitting
from this LPO is predicted for vertical foliation
and horizontal lineation (Fig. 1). The predictions
of Tommasi et al. (1999) for wrench deformation
are shown in Figure 7.
Transcurrent boundaries have a very straightforward kinematic relation with respect to lithospheric coupling. Major strike-slip faults within
the continental lithosphere cut the entire crust
and continue into the mantle lithosphere (Stern
& McBride 1997; Hole et al. 1998). Fast polarization directions are typically subparallel or at a
low angle to major active faults, .including the
San Andreas fault (California; Ozalaybey &
54
B. TIKOFF E T A L .
Fig. 6. Block diagrams of idealized deformation of the lithosphere. Wrench deformation (a) occurs in transcurrent plate
boundaries and results in plane strain fabrics. Oblique divergence results in wrench-dominated transtension (b) or pure
shear-dominated transtension (c), both of which result in constrictional fabrics. Pure shear-dominated transtension
causes horizontal foliations. Pure shear-dominated transpression (d) and wrench-dominated transpression (e) result from
oblique plate convergence. Both are associated with flattening fabrics. Pure shear-dominated transpression should result
in vertical lineations, which cause low amounts of shear-wave splitting in the absence of lateral extrusion. Deformation
that is dominantly divergent or convergent (e.g. pure shear-dominated) will tend towards plane strain fabrics.
M A N T L E - D R I V E N D E F O R M A T I O N OF OROGENS
Fig. 7. Deformation types and the orientation of crystallographic axes of olivine, from predictions of a
polycrystalline plasticity model valid for high-temperature (>900 ~ deformation in the mantle. The deformation
types are shown as blocks on the left side of the diagram. Lower-hemisphere, equal-area stereonets are shown for the
olivine crystallographic axes for 1000 grains contoured at 1% intervals. The black box shows the maximum LPO in
each stereonet. Two predictions of LPO are shown for transpression, distinguishing wrench-dominated (top - long
axis of finite strain ellipsoid is horizontal) and pure shear-dominated (bottom - long axis of finite strain ellipsoid is
vertical). In general, olivine LPO is subparallel to the finite strain axes. Modified from Tommasi et al. (1999).
55
56
B. TIKOFF ET AL.
Savage 1995), E1 Pilar fault (Venezuela; Russo
et aL 1996), Kunlun fault (Tibet; McNamara
et al. 1994) and Altyn Tagh (Tibet; Herquel et al.
1995). Note that, in the case of the San Andreas
fault, only the topmost of two layers of anisotropic
material is oriented at a low angle to the San
Andreas fault (Ozalaybey & Savage 1995). For
the San Andreas and New Zealand cases, the
mantle fabric appears to rotate towards parallelism
with the boundary as one approaches the fault.
This observation suggests that the deformation
gradient is highest under the fault, but that deformation in the mantle occurs broadly under these
plate boundaries.
The same is true for the crustal deformation, as
all of the transcurrent motions of these boundaries
are not accommodated on the major, plateboundary faults. Diffuse crustal deformation
accommodating transcurrent motion was proposed for the San Andreas system by Jamison
(1991) and Teyssier & Tikoff (1998), and is
observed geodetically and structurally for the
South Island of New Zealand (e.g. Beavan &
Haines 2001; Little et aL 2002). Similar patterns
are inferred from major, ancient strike-slip
faults, such as the Great Glen fault (Scotland;
Helffrich 1995) and strike-slip zones in the
Appalachians (Barruol et al. 1997). As addressed
earlier, this fabric survives, and potentially
even controls, subsequent continental break-up
(Vauchez et al. 1997; Tommasi & Vauchez 2001).
D i v e r g e n c e (Fig. 6b, c)
Oblique divergence will typically result in transtensional deformation, provided that a straight
margin exists with a consistent oblique divergence
movement direction. Transtension deformations
typically lead to constrictional fabrics, although
the exact nature of the fabric depends on the
angle of oblique divergence and the amount of
deformation (Fossen & Tikoff 1993; Fossen
et al. 1994). Wrench-dominated transtension
occurs for angles of oblique divergence
less than 20 ~ Except settings that are within a
few degrees from pure transcurrent motion,
wrench-dominated transtension will result in
very constrictional deformation. For pure sheardominated transtension and higher angles of
oblique divergence, the deformation will start
to approach plane strain controlled by the divergent component of motion. The lineation is parallel to the oblique movement direction and not
the shear zone boundary. Foliation is vertical
for low finite strain and low angles of oblique
divergence. For any oblique divergence direction
greater than 20 ~ the foliation is always horizontal (e.g. Fossen et al. 1994). The LPO of olivine
tracks these directions well (Fig. 7), as determined by a polycrystalline plasticity model
valid for high-temperature (>900 ~
mantle
deformation (Tommasi et al. 1999). In general,
less shear-wave splitting is expected for obliquely divergent settings, because of the horizontal foliation (Fig. 1).
Divergence in continental lithosphere results
in extensional or transtensional deformation,
depending on the angle of oblique divergence.
Mantle deformation shows characteristic fabrics
with the fast split wave polarized oblique to the
regional structural trend. The Rio Grande
(Sandvol et al. 1992) and East Africa (Kenya;
Gao et al. 1997) rifts both display this pattern.
Lithospheric thinning may follow the structural
expression of the rift, as demonstrated in Kenya
(Achauer et al. 1994). Upper crustal deformation
responds to this deformation by undergoing a
combination of strike-slip and normal faulting
(e.g. Doser & Yarwood 1991). The similarity
of upper crustal deformation patterns is closely
consistent with experimental models of bottomcontrolled deformation (Withjack & Jamison
1986).
Polycrystalline plasticity models predict that
olivine LPO developed in response to a transtensional deformation will be oblique to the rift
trend, giving rise to a seismic anisotropy signal
characterized by fast split shear wave polarized
in a direction oblique to the rift elongation
(Tommasi et al. 1999). This is consistent with
the orientation of the shear-wave splitting data
observed in the continental rifts. In our model
(Fig. 6b, c), upwelling and outward flowing asthenosphere provides the basal drag to cause the
formation of fabrics in the lithospheric mantle.
As regions experiencing oblique divergence typically have high heat flow, such as the Basin and
Range region of the western United States, it is
assumed that only the uppermost part of the
mantle forms an LPO. This occurs because
high temperatures tend to favour diffusional
creep, which does not produce an LPO.
We have not included passive rifting and/or
orogenic collapse in the above model, despite
the fact that these processes undoubtedly exist.
Thickening of the crust and lithosphere inevitably introduces thermo-mechanical effects, such
as thermal relaxation and time-dependent
heating. These effects would act to decrease the
coupling between the crust and the upper
mantle (Vanderhaeghe & Teyssier 2001). These
extensional systems may undergo top-driven collapse (e.g. Axen et al. 1998), bottom-driven collapse, or some combination, with the variation
occurring on both a spatial and temporal basis
(e.g. Rey et al. 2002; Tikoff et al. 2002). The
MANTLE-DRIVEN DEFORMATION OF OROGENS
connection between these processes and mantle
flow requires more study.
C o n v e r g e n c e (Fig. 6d, e)
Convergence plate boundaries include oceanicoceanic, continental-oceanic, or continentalcontinental settings. We will concentrate on
oceanic-continental and continental-continental
settings. Oblique convergence results in transpressional deformation. Transpression causes
flattening fabrics, which approach plane strain
(pure shear) at very high angles of convergence.
Foliation is always vertical, although lineation
can be horizontal or vertical. Horizontal lineations, which result in strong shear-wave splitting,
occur for an oblique convergence direction less
than 20 ~ and for low strains (e.g. Fossen &
Tikoff 1993; Fossen et al. 1994). The lineation
starts at less than 45 ~ to the shear zone boundary
(plate boundary) and rotates into parallelism with
the shear zone boundary at higher finite strain.
Polycrystalline plasticity models (Tommasi
et al. 1999) show good agreement between
olivine LPO and finite strain axes (Fig. 7). Consequently, strong shear-wave splitting is predicted for the case of horizontal lineations. For
the pure shear end-member, however, the vertical lineations provide very low to no shearwave splitting.
Convergent settings show the largest difference between inferred mantle fabric from seismology and crustal deformation. In most
orogenic zones, reverse faulting and associated
folding show large amounts of crustal contraction perpendicular to the trend of the belt. In
contrast, shear-wave splitting is generally
perpendicular to the maximum shortening direction, for both neotectonic (Tien Shan, Wolfe &
Vernon 1998) and ancient examples (Hercynian,
Bormann et al. 1993; Ribeira belt, Heinz et al.
2003). Note, however, that some measurements
are parallel to the maximum shortening direction in the above cases. From these data,
mantle deformation is inferred to extend
material perpendicular to the contraction direction and parallel to the trend of the belt. As indicared by Tommasi et al. (1999), this pattern
cannot result from a pure shear deformation in
the mantle in the plane of the shortening direction. Pure shear would create a vertical long
axis which is incompatible with the large
( > 1 s) shear-wave splitting observed in these
areas.
Ways of interpreting the data include:
(1) simple shear deformation; (2) pure shear
deformation acting in a horizontal plane
(lateral extrusion); or (3) a combination of the
57
two (Fig. 4). Geodetic and structural studies
are not consistent with large amounts of
strike-slip movement in many foreland fold
and thrust belts. In the Himalayas, shear-wave
splitting increases as one approaches the major
strike-slip faults, including the Kunlun and
Altyn Tagh (Fig. 4). The other possibility is
coaxial deformation, lateral extrusion, which
elongates material parallel to the strike of the
fold and thrust belt. (We use 'lateral extrusion'
to indicate a pure shear that both elongates
and shortens material in a horizontal plane,
although other kinematic possibilities for
lateral extrusion are possible.) Lateral extrusion
would explain the nearly exact perpendicular
relation between shear-wave splitting and crustal
contraction.
Lateral extrusion requires some level of partial
detachment (i.e. clutch tectonics). The shearwave splitting data indicate that the upper
mantle is undergoing contractional flow below
the thrust belt, in a similar direction to the
crustal deformation. However, because of the
different boundary conditions, the direction of
extension is different for the crust and mantle.
Because of the free surface of the Earth, the
extension direction is upward and crustal deformation is characterized as reverse faulting. In
the mantle, material cannot move downward
easily, as it must displace other material.
Therefore, the mantle moves laterally. Alternatively, both simple shear and a component of
lateral extrusion act together (plane transpression).
If extrusion tectonics is applicable to the
Himalayan collision (e.g. Tapponnier et al.
2002), it is worth considering the role of mantle
flow, in a bottom-driven system, and its effect
on crustal deformation. The crustal deformation
is only partially coupled from this extrusion,
and accommodates more contractional deformation. A normal fault system occurs over
most of the Tibetan plateau, which strikes
north-south (parallel to the contraction direction) (Armijo et al. 1986; Yin et al. 1999).
These normal faults are potentially recording a
component of the extension deformation
imposed by the extruding mantle flow in addition
to lower crustal flow (Fig. 4). However, it is
useful to remember that strain is a result of displacement gradient. This requires that the displacement from mantle extrusion increases from
west to east (or, possibly, although less likely,
from east to west) across Tibet. While the
above discussion does not prove that the Himalayas are a bottom-driven system, it does indicate
that a bottom-driven system is kinematically
viable.
58
B. T1KOFF ETAL.
Strike-slip partitioning and homogeneous
mantle deformation
Strike-slip faults, which parallel the trend of the
orogen, occur in most mountain belts. Although
described originally for obliquely convergent
arc settings (e.g. Fitch 1972), strike-slip faulting
also occurs in continental collisions, such as the
Himalayas, with dominantly convergent motion
(e.g. strike-slip partitioning; Teyssier et al. 1995).
If the upper crust is coupled to the mantle, the
zones of strike-slip motion in the upper crust
must represent zones of strong displacement gradients in the mantle. The physical experiments of
Richard & Cobbold (1990) are a good analogy
for this deformation. In these experiments, a
layer of sand overlays a viscous silicone layer
and the deformation was controlled by the
bottom of the experiment. The sand developed
a strike-slip fault immediately above maximum
gradient of displacement in the underlying silicone. The sand on either side of the fault was
deformed, often with reverse faults, but which
accommodated a component of wrenching.
If the upper mantle partitioned deformation at
a scale smaller than a seismic wave, one would
expect shear zones parallel to the plate margin
accommodating the transcurrent motion and
lithons between the shear zones accommodating
the contraction (Fig. 8). As the shear waves
would preferentially travel through the larger
lithons, this would cause fabric parallel to the
major faults. The absence of this pattern (e.g.
South Island of New Zealand) indicates that the
mantle does not partition deformation into
discrete high-strain zones. Rather, the correspondence of crustal deformation and shear-wave
splitting requires that the mantle deforms in a
relatively homogeneous manner.
Strike-slip partitioning, and most other crustal
features in orogenic zones, may thus depend
more on the mantle flow than local strength heterogeneities in the crust. The kinematic connection between movement in the mantle and crust
may occur despite the major difference in the
style of deformation: the upper crust deforms
by faulting while the mantle deforms by crystalplastic deformation. Although plate-margin-parallel, strike-slip faults are locally necessary to
accommodate mantle flow, and they are not
capable of accommodating contraction or extension. Consequently, large regions of contractional or extensional deformation, exhibiting
reverse or normal faults, must accommodate the
remaining part of the imposed mantle displacement. Different types of faults (e.g. strike-slip
and reverse) must coexist to accommodate
brittle deformation driven by movement of a
ductile substrate. Consequently, the concept of
a single regional stress direction that typifies an
orogen is not compatible with a bottom-driven
system. The simple relation between stress and
flow may occur exclusively in the mantle, while
the crust deforms passively in response to an
imposed basally driven displacement field. To
use the vernacular, the interesting physics may
all be in the mantle while the crust tides along.
Absolute plate motion (APM) v. relative
plate motion (RPM)
A potentially productive approach is to study the
difference between absolute plate motion (APM)
Fig. 8. (a) Orientation of finite strain for
homogeneous and partitioned transpression, a is the
angle of oblique convergence. The magnitude of finite
strain is proportional to the amount of LPO and,
therefore, to the amount of shear-wave splitting. (b)
Partitioned delbrmation at any scale less than a seismic
wave (e.g. shear bands) will tend to reduce the amount
of splitting and the angle between the maximum
splitting and the boundary. The correspondence
between crustal deformation and shear-wave splitting
(e.g. Little et al. 2002) indicates that this scale of
deformation partitioning does not occur in the mantle.
MANTLE-DRIVEN DEFORMATION OF OROGENS
and relative plate motion (RPM) for plate
margins with large transcurrent components of
motion. The APM is fixed with respect to hotspots. The San Andreas system in central California
and E1 Pilar system in Trinidad are a particularly
useful comparison (Fig. 9), because of their
difference in APM. Both margins record largely
transcurrent motion and similar amounts of displacement (e.g. Crowell, 1966; Speed 1985).
The San Andreas fault, defined as the plate
boundary between the North American and the
Pacific plates, moves southwestward in an
absolute reference frame. Consequently, the
San Andreas fault moves approximately perpendicular to its strike relative to the mantle (Figs 3 &
9). Essentially, the plate boundary between the
North American and Pacific plate moves
westward, overrunning its old position in the
asthenosphere. In contrast, the plate boundary
between the South American and Caribbean
plates, reflected by the E1 Pilaf fault of Venezuela and its equivalent in Trinidad, is fixed in
an absolute reference frame (Figs 3 & 8).
There is a discernible effect of APM on the
deformation fabric, as reflected in the shearwave splitting observed in both places. The San
Andreas system has low splitting values, with
delay travel times only up to 1.25 s (Hartog &
Schwartz 2001). In fact, models for interpreting
shear-wave splitting in this area require a twolayer anisotropy model, in which only the
upper layer reflects a fabric at a low angle to
the San Andreas fault (e.g. Ozalaybey &
Savage 1995; Hartog & Schwartz 2001). In
general, within California, there is a gradual
San Andreas type
"~
rotation in the polarization of fast split shear
waves from N E - S W in easternmost California
to east-west in central California. This pattern
can be interpreted as a large-scale drag in the
mantle, reflecting a change from olivine LPO
oriented parallel to the APM to that reflecting
the San Andreas fault.
Trinidad/Venezuela has relatively large splitting (Russo et al. 1996), with delay travel times
of ~ 2 s. The large splitting indicates that both
the lithospheric mantle and part of the asthenospheric mantle are pervasively deformed. The
data also show alignment subparallel ( ~ 5 - 8 ~
obliquity) to the Caribbean/South American
plate boundary, unlike other transcurrent plate
boundaries (San Andreas, New Zealand), where
the splitting is at a higher angle to the major
transcurrent faults. Because the Caribbean/
South American plate boundary is fixed in an
absolute reference frame, we interpret the nearparallelism and high delay times to reflect the
higher finite strains recorded along this
boundary.
Therefore, it appears that relative motion is not
the only factor in the development of mantle
fabric. Rather, there are also important structural
implications and differences in mantle fabric
resulting from the absolute plate motion. For
instance, the Caribbean/South American plate
boundary is a relatively narrow and welldefined boundary. There are essentially two
plates that have relative movement and a
~ 1 0 0 k m wide deformation zone between
them. The strong fabrics and narrow deformational zone may result from the fixed position
Trinidad type
RPM:
59
lut
relative plate motion
""
Fig. 9. The role of absolute plate motion in mantle deformation, as shown by a comparison of the San Andreas and
Trinidad systems. Both systems have approximately transcurrent relative motion. The San Andreas system moves
WSW, with respect to a fixed mantle framework. Consequently, the lithospheric wrenching deformation (e.g. fault
system) is moving with respect to the asthenosphere. This may result in low amounts of splitting and very
heterogeneous deformation. In contrast, crustal movement in Trinidad is parallel to the absolute plate motion and the
plate boundary does not move in a fixed mantle framework. The parallelism of shear-wave splitting and the
narrowness (~ 100 km) of the deformation zone are a possible result of the lithospheric boundary being coincident
with an asthenospheric boundary.
60
B. TIKOFF ETAL.
of the plate boundary in the mantle. In contrast,
the P a c i f i c / N o r t h A m e r i c a plate b o u n d a r y in
central California is wide and complicated.
Instead of two plates, there is also the Sierra
Nevada-Great
Valley microplate that is
m o v i n g to the northwest, although at a decreased
rate. This complexity m a y result from the
southwestward m o v e m e n t of the San Andreas
fault system, in an absolute motion framework.
Conclusions
The similarity of deformation patterns in the
upper crust (structural reconstruction, geodetic
data, etc.) and upper mantle (shear-wave splitting, azimuthal anisotropy of Pn waves, etc.)
indicate that these lithospheric layers are, at
least, partially coupled. Crustal rocks in these
deforming zones are hypothesized to be m o v i n g
on a thin and flowing mantle, inferred from
seismic attenuation studies. This deforming
mantle controls orogenesis. A p p l y i n g the clutch
m o d e l - partial attachment b e t w e e n lithospheric
layers - allows us to address three-dimensional
models of convergent, transcurrent and divergent
plate margins, for both crustal and mantle deformation. The history of the plate margin and the
absolute plate motion are critical, as both influence the mantle fabric and overlying crustal
structure s.
BT and CT would like to thank the public library in Black
River Falls, Wisconsin, for providing a meeting room, in
which the first version of this paper was outlined. The
authors would like to thank Shun Karato for nmnerous
conversations about olivine deformation. BT wishes to
thank Scott Giorgis and Eric Horsman for work on the
manuscript that allowed it to be submitted at all.
References
ABRAMSON, E.H., BROWN, J.M., SLUTSKY, L.J. &
ZAUG, J. 1997. The elastic constants of San
Carlos Olivine to 17 GPa. Journal of Geophysical
Research, 102, 12253-12263.
ACHAUER, U., GLAHN, A., et al. 1994. New ideas on
the Kenya Rift based on the inversion of the combined dataset of the 1985 and 1989/90 seismic
tomography experiments. In: PRODEHL, C.,
KELLER, G.R. & KHAN, M.A. (eds) Crustal and
upper mantle structure of the Kenya Rift. Tectonophysics, 236, 305-329.
AMATO, A., CHIARABBA, C. & SELVAGGL G. 1997.
Crustal and deep seismicity in Italy (30 years
after). Annali di Geofisica, 40, 981-993.
ARMIJO, R., TAPPONNIER,P., MERCIER, J.-L. & HAN,
T.L. 1986. Quaternary extension in southern Tibet;
field observations and tectonic implications.
Journal of Geophysical Research, 91, 1380313872.
AXEN, G.J., SELVERSTONE, J., BYRNE, T. &
FLETCHER, J.M. 1998. If the strong crust leads,
will the weak crust follow? GSA Today, 8, 1-8.
BA1, Q., MACKWELL, S.J. & KOHLSTEDT, D.L. 1991.
High-temperature creep of olivine single crystals,
1. Mechanical results for buffered samples.
Journal of Geophysical Research, 96, 2441-2463.
BARRUOL, G. & KERN, H. 1996. Seismic anisotropy
and shear-wave splitting in lower-crustal and
upper-mantle rocks from the Ivrea Zone: experimental and calculated data. In: MAINPRICE, D. &
VAUCHEZ, A. (eds), Dynamics of the subcontinental mantle; from seismic anisotropy to mountain
building. Physics of the Earth and Planetary
Interiors, 95, 175-194.
BARRUOL, G., HELFFRICH, G. & VAUCHEZ, A. 1997.
Shear wave splitting around the northern Atlantic;
frozen Pangaean lithospheric anisotropy? In:
MAINPRICE, D., BOUDIER, F. & BOUCHEZ, J.L.
(eds) The Adolphe Nicolas volume. Tectonophysics, 279, 135-148.
BEAVAN, J. & HAINES, J. 2001. Contemporary horizontal velocity and strain rate fields of the
Pacific-Australian plate boundary zones through
New Zealand. Journal of Geophysical Research,
106, 741-770.
BEN ISMAEL, W. 8,: MAINPRICE, D. 1998. An olivine
fabric database: an overview of upper mantle
fabrics and seismic anisotropy. Tectonophysics,
296, 145-158.
BORMANN, P., BURGHART, P.T., MAKEYEVA, L.I. &
VINNIK, L.P. 1993. Teleseismic shear-wave splitting and deformations in central Europe. Physics
of the Earth and Planetary Interiors, 78, 157-166.
BOSTOCK, M.G. & CASSIDY, J.F. 1995. Upper mantle
stratigraphy beneath the southern Slave Craton.
Canadian Journal of Earth Science, 34, 577-587.
BRACE W.F. & KOHESTEDT D.L. 1980. Limits on
lithospheric stress imposed by laboratory experiments. In: HANKS, T.C. & RALEIGH, C.B. (eds),
Magnitude of deviatoric stresses in the Earth's
crust and uppermost mantle. Journal of Geophysical Research, 85, 6248-6252.
BURG, J.P., DAVY, P. & MARTINOD, J. 1994. Shortening of analogue models of the continental lithosphere; new hypothesis for the formation of the
Tibetan Plateau. Tectonics, 13, 475-483.
BYSTRIKY, M., KUNZE, K., BURLINI, L. & BURG, J.P.
2000. High shear strain of olivine aggregates: rheological and seismic consequences. Science, 290,
1564-1567.
CALVERT, A., SANDVOL, E., et al. 2000. Geodynamic
evolution of the lithosphere and upper mantle
beneath the Alboran region of the western Mediterranean: constraints from travel time tomography.
Journal of Geophysical Research, 105, 1087110898.
CARLSON, R.L. 1995. A plate cooling model relating
rates of plate motion to the age of the lithosphere
at trenches. Geophysical Research Letters, 22,
1977-1980.
CARTER, N.L. & AVE LALLEMANT, H.G. 1970. High
temperature flow of dunite and peridotite. Geological Society of America Bulletin, 81, 2180-2202.
MANTLE-DRIVEN DEFORMATION OF OROGENS
CHASTEL, Y.B., DAWSON, P.R., WENK, H.R. &
BENNETT, K. 1993. Anisotropic convection with
implications for the upper mantle. Journal of Geophysical Research, 98, 17757-17771.
CHRISTENSEN, J.N., ROSENFELD, J.L. & DEPAOLA,
D.J. 1989. Rates of tectonometamorphic processes
from rubidium and strontium isotopes in garnet.
Science, 244, 1465-1469.
CHRISTENSEN, N.I. 2002. Continental mantle seismic
anisotropy: a new look at the Twin Sisters massif.
Tectonophysics, 355, 163-170.
CHRISTENSEN, N.I. & MOONEY, W.D. 1995. Seismic
velocity structure and composition of the continental crust; a global view. Journal of Geophysical
Research, 100, 9761-9788.
CROWELL, J.C. 1966. Displacements along the San
Andreas fault, California. Annales Academia
Scientiarum Fennicae, Series A3: GeologicaGeographica, 90, 119-125.
DOSER, D.I. & YARWOOD,D.R. 1991. Strike-slip faulting in continental rifts; examples from Sabukia,
East Africa (1928), and other regions. In: Gangi,
A.F. (ed) World rift systems. Tectonophysics,
197, 213-224.
DUNLAP, W.J., HIRTH, G. & TEYSSIER, C. 1997.
Thermomechanical evolution of a ductile duplex.
Tectonics, 16, 983-1000.
ENDERLE, U., MECHIE, J., SOBOLEV, S. & FUCHS, K.
1996. Seismic anisotropy within the uppermost
mantle of southern Germany. Geophysical
Journal International, 12S, 747-767.
FITCH, T.J. 1972. Plate convergence, transcurrent
faults, and internal deformation adjacent to Southeast Asia and the Western Pacific. Journal of
Geophysical Research, 77, 4432-4460.
FOSSEN, H. & TIKOFF, B. 1993. The deformation
matrix for simultaneous simple shearing, pure
shearing and volume change, and its application
to transpression-transtension
tectonics. In:
CASEY, M., DIETRICH, D., FORD, M., WATKINSON,
J. & HUDLESTON, P.J. (eds) The geometry of naturally deformed rocks. Journal of Structural
Geology, 15, 413-422.
FOSSEN, H., TIKOEF, B. & TEYSSIER, C. 1994. Strain
modeling of transpression and transtensional deformation. In: KRILL, A.G. & LIPPARD, S.J. (eds) Strike-slip tectonics: Proceedings of the 10th TSGS
conference. Norsk Geologisk Tidsskrif, 74, 134-145.
GAO, S., DAVIS, P.M., et al. 1997. SKS splitting
beneath continental rift zones. Journal of Geophysical Research, 102, 22781-22797.
GIORGIS, S., MARKLEY,M., and TIKOFF, B. 2004. Vertical-axis rotation of rigid crustal blocks driven by
mantle flow. In: GROCOT, J., MCCAFFREY, K.J.W.,
TAYLOR, G. & TIKOFF, B. (eds) Vertical Coupling
and Decoupling in the Lithosphere. Geological
Society, London, Special Publications, 227, 83-99.
GOSCOMBE, B. 1992. High-grade reworking of central
Australian granulites: Part 1, Structural evolution.
Tectonophysics, 204, 361-399.
HARTOG, R. & SCHWARTZ,S.Y. 2001. Depth-dependent
mantle anisotropy below the San Andreas Fault
system; apparent splitting parameters and wave-
61
forms. Journal of Geophysical Research, 160,
4155-4167.
HATZFELD, D., KARAKOSTAS, V., et al. 2000. Microseismicity and faulting geometry in the Gulf of
Corinth (Greece). Geophysical Journal International, 141, 438-456.
HEARN T.M. 1996. Anisotropic Pn tomography in the
Western United States. Journal of Geophysical
Research, 101, 8403-8414.
HEARN T.M. 1999. Uppermost mantle velocities and
anisotropy beneath Europe. Journal of Geophysical
Research, 104, 15123-15139.
HEINTZ, M., VAUCHEZ, A., ASSUMPCAO, G.,
BARRUOL, G. & E~YOIO-SILVA, M. 2003. Shear
wave spilling in SE Brazil: an effect of active on
fossil upper mantle flow, or both? Earth and Planetary Science Letters, 211, 79-95.
HELFFRICH, G. 1995. Lithospheric deformation
inferred from teleseismic shear wave splitting
observations in the UK. Journal of Geophysical
Research, 100, 18195-18204.
HERQUEL, G., WITTLINGER, G. & GUILBERT, J. 1995.
Anisotropy and crustal thickness of NorthernTibet: new constraints for tectonic modelling. Geophysical Research Letters, 22, 1925-1928.
HESS, H.H. 1964. Seismic anisotropy of the uppermost
mantle under oceans. Nature, 203, 629-631.
HOLE J.A., BEAUDOIN, B.C. & HENSTOCK, T.J. 1998.
Wide-angle seismic constraints on the evolution of
the deep San Andreas plate boundary by Mendocino triple junction migration. Tectonics, 17,
802-818.
HOLT, W.E. 2000. Correlated crust and mantle strain
fields in Tibet. Geology, 28, 67-70.
IMBER, J., HOLDSWORTH,R.E. & BUTLER, C.A. 2001.
A reappraisal of the Sibson-Scholz fault zone
model: the nature of the frictional to viscous
('brittle-ductile') transition along a long-lived,
crustal-scale fault, Outer Hebrides, Scotland. Tectonics, 20, 601-624.
JAMES, D.E. & ASSUMPCAO, M. 1996. Tectonic
implications of S-wave anisotropy beneath SE
Brazil. Geophysical Journal International, 126,
1-10.
JAMISON, W.R. 1991. Kinematics of compressional
fold development in convergent wrench terranes.
Tectonophysics, 190, 209-232.
JORDAN, T.H. 1975. The continental tectosphere.
Reviews of Geophysics and Space Physics, 13,
1-12.
KAMINSKI, E. & RIBE, N.M. 2001. A kinematic model
for recrystallization and texture development in
olivine polycrystals. Earth and Planetary Science
Letters, 189, 253-267.
KENNEDY, L.A., RUSSEL, J.K. & KOPYLOVA, M.G.
2002. Mantle shear zones revisited: the connection
between the cratons and mantle dynamics.
Geology, 30, 419-422.
KLOSKO, E., Wu, F., et al. 1999. Upper mantle
flow in the New Zealand region from seismic
anisotropy. Geophysical Research Letters, 26,
1497-1500.
KOHLSTEDT, D.L., EVANS, B. & MACKWELL, S.J.
1995. Strength of the lithosphere: constraints
62
B. TIKOFF ETAL.
imposed by laboratory experiments. Journal of
Geophysical Research, 100, 17587-17602.
LAVE, J., AVOUAC,J.P., LACASSIN,R., TAPPONNIER,P.
& MONTAGNER, J.P. 1996. Seismic anisotropy
beneath Tibet; evidence for eastward extrusion of
the Tibetan lithosphere? Earth and Planetary
Science Letters, 140, 83-96.
LAY, T. & WALLACE, T.C. 1995. Modern Global
Seismology. Academic Press, San Diego, CA.
LIN, S. & WILLIAMS, P.F. 1992. The geometrical
relationship between the stretching lineation and
the movement directions of shear zones. Journal
of Structural Geology, 14, 491-497.
LITTLE, T.A., SAVAGE, M.K. & TIKOFF, B. 2002.
Relationship betweem crustal finite strain
and seismic anisotropy in the mantle, PacificAustralia plate boundary zone, South Island,
New Zealand. Geophysical Journal International,
151, 160-169.
MACKWELL,S.J., KOHLSTEDT,D.L. & PATERSON,M.S.
1985. The role of water in the deformation of
olivine single crystals. Journal of Geophysical
Research, 90, 11319-11333.
MARGHERIT1,L., NOSTRO, C., AMATO,A. & Cocco, M.
1996. Seismic anisotropy; an original tool to understand the geodynamic evolution of the Italian
Peninsula. In: MANTOVANI, E. (ed.) Geodynamics
of the Mediterranean region and its implications
for seismic and volcanic risk. Annali di Geofisica,
40, 759-769.
MARSON-PIDGEON, K.A. & SAVAGE, M.K. 1997.
Frequency-dependent anisotropy in Wellington,
New Zealand. Geophysical Research Letters', 24,
3297-3300.
MCKENZIE, D. & JACKSON, J. 1983. The relationship
between strain rates, crustal thickening, palaeomagnetism, finite strain and fault movements
within a deforming zone. Earth and Planetary
Science Letters, 65, 182-202.
MCNAMARA, D.E., OWENS, T.J., SILVER, P.G. &
Wu, F.T. 1994. Shear wave anisotropy beneath
the Tibetan Plateau. Journal of Geophysical
Research, 99, 655-665.
MELE, G., ROVELLI,A., SEBER, D., I-lEARN, T.M. &
BARAZANG1, M. 1998. Compressional velocity
structure and anisotropy in the uppermost mantle
beneath Italy and surrounding regions. Journal of
Geophysical Research, 103, 12529-12543.
MOCHIZUKI, K., KASAHARA, J., SATO, T.,
S~INOHARA, M. & HIRATA, N. 1997. Seismic
refraction tomography and its application in an
experiment in the Tsushima Basin, Japan Sea.
Geophysical Exploration, 50, 179-195.
MOLNAR, P. 1992. Brace-Goetze strength-profiles,
the partitioning of strike-slip and thrust faulting at
zones of oblique convergence, and the stressheat flow paradox of the San Andreas fault. In:
EVANS, B. & WONG, T.-F. (eds) Fault Mechanics
and Transport Properties of Rocks. Academic
Press, New York, 435-459.
MOLNAR, P., ANDERSON, H.J., et al. 1999. Continuous
deformation versus faulting through the continental
lithosphere of New Zealand. Science, 286, 516519.
MONTAGNER, J.-P. & TANIMOTO, T. 1990. Global anisotropy in the upper mantle inferred from the regionalization of phase velocities. Journal of
Geophysical Research, 95, 4797-4819.
MONTAGNER, J.-P. & TANIMOTO, T. 1991. Global
upper mantle tomography of seismic velocities
and anisotropies. Journal of Geophysical Research,
96, 20337 -20351.
MOORE, M., ENGLAND, P. & PARSONS, B. 2002.
Relation between surface velocity field and shearwave splitting in the South Island of New
Zealand. Journal of Geophysical Research, 107,
doi: 10.1029/2000JB000093.
NISHIMURA, C.E. & FORSYTH, D.W. 1988. Rayleigh
wave phase velocities in the Pacific with implications for azimuthal anisotropy and lateral heterogeneities. Journal of Geophysical Research, 94,
479-50I.
NISHIMURA, C.E. & FORSYTH, D.W. 1989. The anisotropic structure of the upper mantle in the Pacific.
Geophysical Journal of the RAS, DGG, and EGS,
96, 2O3-229.
OKAL, E.A. & TALANDIER, J. 1997. T waves from the
great 1994 Bolivian deep earthquake in relation to
channeling of S wave energy up the slab. Journal of
Geophysical Research, 102, 27421-27437.
OLDOW, J.S., BALLY, A.W. & AVE LALLEMANT, H.G.
1990. Transpression, orogenic float, and lithospheric balance. Geology, 18, 991-994.
[~)ZALAYBEY, S. ~ SAVAGE, M.K. 1995. Shear-wave
splitting beneath western United States in relation
to plate tectonics. Journal of Geophysical
Research, 100, 18135 - 18149.
PHAKEY, P., DOLL1NGER, G. & CHRISTIE, J. 1972.
Transmission electron microscopy of experimentally deformed olivine crystals. In: Flow and
fracture of rocks. Geophysical Monograph, 16,
117-138.
PLOMEROVA, J., BABUSKA, V. & HORALEK, J. 1998.
Seismic anisotropy and velocity variations in the
mantle beneath the Saxothuringicum-Moldanubicum contact in Central Europe. In: PLOMEROVA,
J., LIEBERMANN,R.C. & BABUSKA, V. (eds) Geodynamics of lithosphere and Earth's mantle:
seismic anisotropy as a record of the past and
present dynamic processes. Pure and Applied Geophysics, 151, 365-394.
PLOMEROVA, J., KOUBA, D. & BABUSKA, V. 2002.
Mapping the lithosphere-asthenosphere boundary
through changes in surface-wave anisotropy.
Tectonophysics, 358, 175-186.
RAITT, R.W., SHOR, G.G., FRANCIS, T.J.G. & MORRIS,
G.B. 1969. Anisotropy of the Pacific upper mantle.
Journal of Geophysical Research, 74, 3095-3109.
REY, P., VANDERHAEGHE, O. ~:; TEYSS1ER, C. 2002.
Gravitational collapse of continental crust: definitions, regimes, mechanisms and modes. Tectonophysics, 342, 435-449.
RIBE, N.M. & Yu, Y. 1991. A theory for plastic deformation and textural evolution of olivine polycrystals. Journal of Geophysical Research, 96,
8325-8335.
RICHARD, P. & COBBOLD, P. 1990. Experimental
insights into partitioning fault motions in continen-
MANTLE-DRIVEN DEFORMATION OF OROGENS
tal convergent wrench zones. In: BOCCALETTI,
M. & NUR, A. (eds) Active and recent strike-slip
tectonics; selected papers from an international
workshop. Annales Tectonicae, 4, 35-44.
ROTH, E.G., WIENS-DOUGLAS, A., DORMAN, L.M.,
HILDEBRAND, J. & WEBB, S.C. 1999. Seismic
attenuation tomography of the Tonga-Fiji region
using phase pair methods. Journal of Geophysical
Research, 104, 4795-4809.
ROYDEN, L.H., BURCHFIEL, B.C., KING, R.W.,
WANG, E., CHEN, Z., SHEN, F. & LIU, Y. 1997.
Surface deformation and lower crustal flow in
eastern Tibet. Science, 276, 788-790.
Russo, R.M. & OKAL, E.A. 1998. Shear wave splitting
and upper mantle deformation in French Polynesia:
evidence from small-scale heterogeneity related to
the Society Hotspot. Journal of Geophysical
Research, 103, 15089-15107.
Russo, R.M., SILVER,P.G., FRANKE,M., AMBEH,W.G.
& JAMES, D.E. 1996. Shear-wave splitting in
northeast Venezuela, Trinidad, and the eastern
Caribbean. Physics of the Earth and Planetary
Interiors, 95, 251-275.
SANDVOL, E., NI, J., OZALAYBEY, S. & SCHLUE, J.
1992. Shear-wave splitting in the Rio Grande
Rift. Geophysical Research Letters, 19, 23372340.
SAVAGE, M.K. 1999. Seismic anisotropy and mantle
deformation; what have we learned from shear
wave splitting? Reviews of Geophysics, 37,
65-106.
SAVAGE, M.K. & SILVER, P.G. 1993. Mantle deformation and tectonics; constraints from seismic
anisotropy in the western United States. In:
MAINPRICE, D., VAUCHEZ, A. & MONTAGNER,
J.P. (eds) Dynamics of the subcontinental mantle:
from seismic anisotropy to mountain building.
Physics of the Earth and Planetary Interiors, 78,
207-227.
SCHOLZ, C.H. 1988. The brittle-ductile transition and
the depth of seismic faulting. Geologisches
Rundschau, 77, 319-328.
SIBSON. R.H. 1977. Fault rocks and fault mechanisms.
Journal of the Geological Society of London, 133,
191-213.
SILVER, P.G. 1996. Seismic anisotropy beneath the
continents: probing the depths of geology. Annual
Review of Earth and Planetary Sciences, 24,
385-432.
SILVER, P.G. & CHAN, W.W. 1991. Shear wave splitting and subcontinental mantle deformation.
Journal of Geophysical Research, 96, 1642916454.
SILVER, P.G. & HOLT, W.E. 2002. The mantle flow
field beneath western North America. Science,
295, 1054-1058.
SPEED, R.C. 1985. Cenozoic collision of the Lesser
Antilles Arc and continental South America and
the origin of the E1 Pilar fault. Tectonics, 4, 41-69.
STERN, T.A. & MCBRIDE, J.H. 1997. Seismic exploration of continental strike-slip zones. In: KLEMPERER, S.L. & MOONEY, W.D. (eds) Deep
seismic profiling of the continents: I, General
63
results and new methods. Tectonophysics, 286,
63 -78.
TAPPONNIER, P., ZHIQIN, X., ROGER, F., MEYER, B.,
ARNAUD, N., WITTLINGER, G. & JINGSUI, Y.
2002. Oblique stepwise rise and growth of the
Tibetan Plateau. Science, 294, 1671 - 1677.
TE'CSSIER, C. & TIKOFE, B. 1998. Strike-slip partitioned transpression of the San Andreas fault
system: a lithospheric scale approach. In:
HOLDSWORTH, R.E., STRACHAN, R.A. & DEWEY,
J.F. Continental Transpression and Transtension
Tectonics. (eds) Geological Society, London,
Special Publications, 135, 143-158.
TEYSSIER, C., TIKOFF, B. & MARKLEY, M. 1995.
Oblique plate motion and continental tectonics.
Geology, 23, 447-450.
TIKOeF, B. & GREENE, D. 1997. Stretching lineations
in transpressional shear zones; an example from
the Sierra Nevada Batholith, California. Journal
of Structural Geology, 19, 29-39.
TIKOFF, B., TEYSS1ER, C. & WATERS, C.L. 2002.
Clutch tectonics and partial attachment of lithospheric layers. LUG Special Publication Series,
1, 93-117.
TOMMASI, A. & VAUCHEZ, A. 2001. Continental
rifting parallel to ancient collisional belts: an
effect of the mechanical anisotropy of the lithospheric mantle. Earth and Planetary Science
Letters, 185, 199-210.
TOMMASI, A., TIKOFF, B. & VAUCHEZ,A. 1999. Upper
mantle tectonics: three-dimensional deformation,
olivine crystallographic fabrics and seismic properties. Earth and Planetary Science Letters, 168,
173-186.
TOMMASI, A., MAINPRICE, D., CANOVA, G. &
CHASTEL, Y. 2000. Viscoplastic self-consistent
and equilibrium-based modeling of olivine lattice
preferred orientations: implications for the upper
mantle seismic anisotropy. Journal of Geophysical
Research, 105, 7893-7908.
VANDERHAEGHE, 0. & TEYSSIER, C. 2001. Crustal
scale rheological transitions during late-orogenic
collapse. Tectonophysics, 335, 211-228.
VAN DER LEE, S. & NOLET, G. 1997. Seismic image of
the subducted trailing fragments of the Farallon
Plate. Nature, 386, 266-269.
VAUCHEZ, A. & GARRIDO, C.J. 2001. Seismic properties of an asthenospherized lithospheric mantle:
constraints from lattice preferred orientations in
peridotite from the Ronda Massif. Earth and Planetary Science Letters, 192, 235-249.
VAUCHEZ, A., TOMMASI, A. & EGYDIO-SANTOS,
M. 1994. Self-indentation of a heterogeneous continental lithosphere. Geology, 22, 967-970.
VAUCHEZ, A., BARRUOL, G. & TOMMASI, A. 1997.
Why do continents break-up parallel to ancient orogenic belts? Terra Nova, 9, 62-66.
VINNIK, L.P., FARRA, V. & ROMANOWICZ, B. 1989a.
Azimuthal anisotropy in the Earth from observations of SKS at GEOSCOPE and NARS broadband stations. Bulletin of the Seismological
Society of America, 79, 1542-1558.
VINN1K,L.P., KIND, R., KOSAREV,G.L. & MAKEYEVA,
L.I. 1989b. Azimuthal anisotropy in the lithosphere
64
B. TIKOFF ET AL.
from observations of long-period S-waves. Geophysics Journal International, 99, 549-559.
VINNIK, L.P., KRISHNA,V.G., KaND, R., BORMANN,P.
& STAMMLER,K. 1994. Shear wave splitting in the
records of the German Regional Seismic Network.
Geophysical Research Letters, 21,457-460.
WENK, H.R. & TOME, C.N. 1999. Modeling dynamic
recrystallization of olivine aggregates deformed
in simple shear. Journal of Geophysical Research,
104, 25513-25527.
WENK, H.R., BENNETT, K., CANOVA, G.R. &
MOLINAm, A. 1991. Modelling plastic deformation
of peridotite with the self-consistent theory.
Journal of Geophysical Research, 96, 83378349.
WITHJACK, M.O. & JAMISON, W.R. 1986. Deformation produced by oblique rifting. Tectonophysics, 126, 99-124.
WOLFE, C.J. & SILVER,P.G. 1998. Seismic anisotropy
of oceanic upper mantle; shear wave splitting
methodologies and observations. Journal of Geophysical Research, 103, 749-771.
WOLFE, C.J. & VERNON,F.L. 1998. Shear-wave splitting at central Tien Shan: evidence for rapid variation of anisotropic patterns. Geophysical
Research Letters, 25, 1217-1220.
YIN, A., KAPP, P.A., et al. 1999. Significant late
Neogene east-west extension in northern Tibet.
Geology, 27, 787-790.
ZHANG, S. & KARATO, S. 1995. Lattice preferred
orientation of olivine aggregates in simple shear.
Nature, 375, 774-777.
ZHANG, S., KARATO, S., FITZGERALD,J., FAUL, U. &
ZHOU, Y. 2000. Simple shear deformation of
olivine aggregates. Tectonophysics, 316, 133-152.